www.geologicacarpathica.sk
GEOLOGICA CARPATHICA, AUGUST 2010, 61, 4, 327—339 doi: 10.2478/v10096-010-0019-y
Post-magmatic hydrothermal mineralization associated with
Cretaceous picrite (Outer Western Carpathians, Czech
Republic): interaction between host rock and externally
derived fluid
ZDENĚK DOLNÍČEK , TOMÁŠ URUBEK and KAMIL KROPÁČ
Department of Geology, Palacký University, Tř. 17. listopadu 12, 771 46 Olomouc, Czech Republic; dolnicek@prfnw.upol.cz
(Manuscript received November 20, 2009; accepted in revised form March 11, 2010)
Abstract: Mineralogy, fluid inclusions, C, O, S, Sr isotopes and trace elements have been studied in amygdule and vein
mineralization hosted by the Lower Cretaceous effusive picrite at Hončova hůrka (Silesian Unit, Flysch Belt of the Outer
Western Carpathians). Besides dominating dolomite, magnesite, siderite, quartz, calcite, chlorite, glauconite, fluorite,
barite, pyrite and millerite were also identified. The parent fluids are characterized by low temperatures (< 50—170 °C), low
salinities (0.4 to 3.7 wt. % NaCl eq.), low content of strong REE-complexing ligands,
δ
18
O,
δ
13
C and
δ
34
S ranges of
0/+ 14 ‰ SMOW, 0/—9 ‰ PDB and ~ 0 ‰ CDT, respectively, and initial
87
Sr/
86
Sr ratios much more radiogenic (0.7060 to
0.7068) than those of host picrite (0.7042 and 0.7046). The fluids are interpreted to be predominantly of external origin,
derived from mixing of seawater with diagenetic waters produced by dewatering of clay minerals in the associated flysch
sediments. The isotope and REE signatures indicate interaction of at least a part of fluids with sedimentary carbonates. The
interaction of fluids with host picrite led to strong alteration of rock-forming minerals and leaching of some elements (Mg,
Ni, S, partly also REE) that widely participated during precipitation of vein- and amygdule-hosted mineral phases.
Key words: Outer Western Carpathians, stable isotopes, strontium isotopes, fluid-rock interaction, veins, fluid inclusions,
amygdules, REE, picrite.
Introduction
The mafic and alkaline-to-subalkaline igneous rocks of the
teschenite association are widespread in the area between
Hranice in the Czech Republic and Bialsko-Biała in Poland,
which belongs to the Silesian Unit of the Outer Western Car-
pathians (Fig. 1). The magmatic rocks forming sills, dikes,
submarine extrusions and pillow lavas are classified as te-
schenites, picrites, monchiquites and alkaline basalts (Pacák
1926; Smulikowski 1929; Kudělásková 1987; Hovorka &
Spišiak 1988). Although these rocks were the subject of
long-lasting scientific interest (the term “teschenite” was de-
fined precisely in this area in the second half of the 19
th
cen-
tury), many aspects of their origin remain unresolved.
A typical feature of these rocks is the intense hydrothermal
alteration, characterized by pervasive chloritization, serpenti-
nization, carbonatization, silicification and zeolitization of pri-
mary magmatic mineral phases (Pacák 1926; Šmíd 1962;
Trundová 2004). Besides the alteration of rock-forming min-
erals, also amygdule fill and mineral veins developed in mag-
matic rocks at many sites. The hydrothermal assemblages are
formed mainly by carbonates, chlorite, quartz, opal or chalce-
dony, locally also zeolites and/or sulphides occur (Bernard et
al. 1981; Sušeň 2000; Urubek 2009). The source of mineraliz-
ing fluids is still under debate. Considering the geological po-
sition of the magmatic suite, the earlier researchers (Pacák
1926; Šmíd 1962) proposed seawater as the main source of
fluids (cf. also Kiss et al. 2008), which can be possibly com-
bined with magmatic waters remaining after crystallization of
host igneous rocks. However, the hypothesis has not been sup-
ported by relevant data. Recently, Dolníček et al. (2010)
showed that magmatic fluids were responsible only for early
mineralization of titanite, aegirine-augite, annite, strontian ap-
atite and analcime, whereas the younger assemblage of
carbonates, chlorite, quartz and sulphides precipitated from an
externally derived fluid supplied by the surrounding sedimen-
tary sequences.
In this contribution we characterize mineralogically and
genetically the hydrothermal mineralization from well-
known mineralogical site at Hončova hůrka Hill. The
mineralogical study extends early descriptive papers by
Rusek & Valošek (1968) and Kudělásková et al. (1990). Fur-
thermore, our study is supplemented by trace element, fluid
inclusion, and C, O, S and Sr isotope data enabling a more
complex evaluation of the origin of the mineralization.
Geological setting
The study area is situated in the Outer Western Carpathians,
an external part of the Alpine-type orogenic belt of the
Western Carpathians (Plašienka et al. 1997). The Externides
consist of Middle-to-Upper Miocene Molasse Basin of the
Carpathian Foredeep, and Upper Jurassic-to-Early Miocene
fold-and-thrust Flysch Belt, representing the Tertiary accre-
tionary wedge of the Carpathian orogen. In the NE part of
the Czech Republic, the Flysch Belt is subdivided into the
Subsilesian, Silesian and Magura Nappes, listed from tectonic
328
DOLNÍČEK, URUBEK and KROPÁČ
foot-wall to hanging-wall (Fig. 1). The studied locality is situ-
ated in the Silesian Unit consisting mainly of Upper Jurassic-
to-Upper Oligocene marine sediments (Eliáš 1970). The basal
carbonate-rich rocks (Těšín Limestone of the Late Jurassic-
Early Cretaceous age) are followed by rhythmic and cyclic
flysch sedimentation of claystones, sandstones, and rare
conglomerates of Early Cretaceous to Late Oligocene age.
Locally, siderite horizons are present in claystone-rich cycles
(e.g. Hradiště Formation, Veřovice Formation), layers of
dark organic-rich biogenic silicite (Menilite Formation) and
bodies of teschenite rocks association (Hradiště Formation).
The whole-rock sequence was folded and thrusted towards
the NW (over the SE part of the Bohemian Massif, Fig. 1)
during the Alpine Orogeny in Paleocene—Miocene times.
Based on the field and geochronological evidence (Lucińska-
Anczkiewicz et al. 2002; Grabowski et al. 2003), the intru-
sion of the teschenite rock association was coeval with
sedimentation of the Lower Cretaceous Hradiště Formation.
The mostly mafic quartz-free alkaline-to-subalkaline rocks
are very variable in texture, mineral composition and chem-
istry, which is interpreted as a result of fractional crystalliza-
tion, mixing of magmas of different origins, assimilation of
sedimentary rocks and post-magmatic alterations (Kudělás-
ková 1987; Hovorka & Spišiak 1988; Dostal & Owen 1998;
Wlodyka & Karwowski 2004). In general, the most primitive
rocks are picrites, and the most evolved and fractionated are
teschenites. The rocks have high concentrations of incompati-
ble elements (Ti, P, Zr, LREE) and low contents of compatible
trace elements (Cr, V). The interpretation of their geotectonic
setting is not conclusive. Narębski (1990) and Spišiak &
Hovorka (1997) stated their geochemical affinity to within-
plate basalts generated during the early rifting in the Silesian
Unit. By contrast, Dostal & Owen (1998) interpreted these
rocks (“lamprophyres”) as post-tectonic intrusions bound to
reactivation of deep faults during the Early Cretaceous.
The studied locality (N49° 39.575’, E18° 09.176’, altitude
317 m a.s.l.) is an old quarry situated 2 km north of the town
of Příbor on the western slope of the Hončova hůrka Hill. The
quarry is situated in a nappe relic of the Silesian Unit, which
overlies the Subsilesian Unit. A dominating part of the quarry
is formed by black-green effusive rocks belonging to peridot-
ite picrite, which often exhibits the porphyric and/or amygdal-
oid texture (Kudělásková 1987; Matýsek 1989). The contact
with clastic sediments (dark claystones with interpositions of
sandy limestones) of the Hradiště Formation is exposed in the
S and SE wall, where thermal metamorphism gave rise to grey
calcareous contact hornfels. Throughout the whole quarry, the
picrite is strongly altered. Phenocrysts of primary mafic
minerals, olivine, clinopyroxene, biotite as well as the ground-
mass (taking 25—36, 13—44, 1—6, and 16—35 vol. %, respec-
tively; Kudělásková 1987) are partly or completely changed
into a mixture of chlorite, serpentine and carbonate. Chemical
compositions of three rock samples that differ in degree of
alteration are given in Table 1. The hydration and carbonatiza-
tion effects are marked by increased loss on ignition (up to
17 wt. %), calcium (up to 12 wt. % CaO) and CO
2
(up to
7.8 wt. %). By contrast, the contents of magnesium and iron
are depleted in altered samples.
Both the picrite and sediments are cut by white to pinkish
hydrothermal veinlets < 1 to 30 mm thick and up to several
meters long. Laterally, the thickness is quite constant for in-
dividual veins. A NW-SE striking system of steep (55—85°)
veinlets has been recognized. The tectonic striae have never
been observed in the vein fill nor on the contact of the rock
Fig. 1. Geological position of the Hončova hůrka locality in the Outer Western Carpathian’s flysch nappe system.
329
HYDROTHERMAL MINERALIZATION IN CRETACEOUS PICRITE (WESTERN CARPATHIANS, CZECH REPUBLIC)
and vein. There are no open drusy cavities, all fissures are
completely cemented. Brecciation of the host rocks has
sometimes been observed.
Methods
The electron microprobe analyses of minerals were per-
formed using a Cameca SX-100 microprobe at the Masaryk
University in Brno. For carbonate and silicate minerals, the
acceleration voltage of 15 kV, 20 nA beam current and beam
diameter of 10 µm (carbonates) and 5 µm (phyllosilicates),
respectively, were used. For sulphide minerals, voltage of
25 kV, beam current of 20 nA and beam diameter of 1 µm
were applied. The collected data were converted to wt. % us-
ing the automatic PAP procedure (Pouchou & Pichoir 1985).
Synthetic phases and well-defined minerals have been used
as standards.
Fluid inclusions were investigated by means of petrography
and optical microthermometry in standard doubly polished
wafers and cleavage fragments. Primary (P), primary-second-
ary (PS) and secondary (S) inclusions were distinguished ac-
cording to the criteria given by Roedder (1984) and Shepherd
et al. (1985). Microthermometric parameters were measured
using a Linkam THMSG 600 stage at Palacký University,
Olomouc. The temperature of final homogenization (T
h
),
eutectic temperature (T
e
) and melting temperature of ice
(T
m
ice) were measured. The stage was calibrated with inor-
ganic standards and synthetic fluid inclusions. The reproduc-
ibility is within 0.1 °C for temperatures between —56.6 and
0 °C, and within 1 °C at 374.1 °C.
For bulk chemical analyses, the carbonate samples weigh-
ing between 1 and 2 g were hand-picked under a binocular
microscope and then pulverized in an agate mortar. The host
rock was powdered in an epicyclic mill and reduced in
weight by quartering. Chemical analyses were performed in
the ACME Analytical Laboratories, Vancouver, Canada.
Aliquots for heavy metal analyses were dissolved in hot
(95 °C) aqua regia and analysed using the ICP-ES method.
Other determined elements including refractory metals and
rare earth elements (REE) were analysed by ICP-MS in an-
other sample aliquot decomposed using LiBO
2
fusion fol-
lowed by leaching in diluted (5%) HNO
3
. Reproducibility of
results is within 5—10 % based on repeated analyses. The Ce
and Eu anomalies were calculated using the equations given
by McLennan (1989).
Stable isotope analyses were conducted in the laboratories
of the Czech Geological Survey, Prague, using a Finnigan
MAT 251 mass spectrometer. The conversion of carbonates
to CO
2
was made by reaction with 100% orthophosphoric
acid (McCrea 1950). For sulphur isotope analyses, the sam-
ples were firstly decarbonatized using 10% HCl. Sulphidic
sulphur in the insoluble residuum was converted to sulphate
using HNO
3
-HCl mixture (Jarchovský 1960), and precipitat-
ed as barium sulphate. The SO
2
gas for sulphur isotope anal-
ysis was produced by heating of BaSO
4
with a SiO
2
+ V
2
O
5
mixture (Ueda & Krouse 1987) at 1050 °C. Results of iso-
tope analyses are conventionally expressed in delta (
δ) nota-
tion as per mil (‰) deviation from commonly used standards
(PDB, SMOW, CDT). Uncertainty is better than ± 0.05, ± 0.1,
and ± 0.2 ‰ for C, O and S isotopic composition, respective-
ly. The isotopic composition of the parent fluid was calculat-
ed using the equations published by O’Neil et al. (1969) and
Ohmoto & Goldhaber (1997). When calculating the fluid
δ
13
C value, equations for H
2
CO
3
and HCO
3
—
as dominating
carbon species were used for temperatures above and below
~ 120 °C (cf. Matsuhisa et al. 1985), respectively.
For strontium isotope determinations, the splits of pow-
dered samples used for trace element determinations were
decomposed using HCl (carbonates) or HNO
3
-HCl-HF mix-
ture (rocks). Strontium was isolated on PP columns with
Sr.spec Eichrom resin. The Sr isotopic ratios were measured
using a Finnigan MAT 262 thermal ionization mass spectrom-
eter in dynamic mode using a double Re filament assembly in
the isotope laboratory of the Czech Geological Survey, Pra-
gue. The
87
Sr/
86
Sr ratios were corrected for mass fractionation
to
86
Sr/
88
Sr = 0.1194. External reproducibility was controlled
by repeated analyses of the NBS 987 (
87
Sr/
86
Sr = 0.710247
± 26 (2
σ), n=25) isotopic standard. The decay constant of
1.42*10
—11
yr
—1
was used for subsequent calculations (Steiger
& Jäger 1977).
Results
Mineralogy
Veins
The mineral composition of all veins is very simple, they
are formed only by carbonates. We have not recently ob-
served zeolite minerals (harmotome, heulandite, ferrierite)
described by Kudělásková et al. (1990). The claystone-host-
ed veins are filled by white coarse-crystalline calcite (iso-
metric grains up to 1 cm in size; sample HH-7). By contrast,
the picrite-hosted ~0.5 cm thick veinlets are formed by white
to pinkish fibrous dolomite, with individual fibres arranged
Table 1: Chemical composition of selected samples from Hončova
hůrka. Analysis No. 44 is taken from Kudělásková (1987). LOI –
loss on ignition, F/FM – FeO/(FeO + MgO) weight ratio, n.d. – not
determined.
HH-4 44 HH-1 HH-3
Sample
Picrite Picrite Picrite Dol.
II+magn
P
2
O
5
0.41
0.65
0.81
0.03
SiO
2
39.09
37.10
35.58
0.75
TiO
2
1.87
2.15
2.66
<0.01
Al
2
O
3
8.00
11.60
10.61
<0.01
Cr
2
O
3
0.13
n.d.
0.11
0.02
FeO
tot
11.25
10.08
7.12
5.60
MnO
0.16
0.18
0.11
0.04
MgO
21.00
17.20
9.45
24.55
CaO
8.87
12.20
12.39
19.50
Na
2
O
1.35
0.75
0.30
0.04
K
2
O
0.99
0.55
1.86
<0.01
LOI
5.00
6.50
17.70
46.70
Total
97.70 98.96 97.88 97.20
CO
2
1.28
4.28
7.81
49.03
S
tot
0.07
n.d.
0.02
<0.02
F/FM
0.35
0.37
0.43
0.19
330
DOLNÍČEK, URUBEK and KROPÁČ
perpendicular to the walls of the vein (sample HH-5). No
remnants of host picrite occur within the vein. The WDX
analyses proved chemical homogeneity of the carbonate con-
taining ~ 91, ~ 8.5 and ~ 0.5 mol % of dolomite, ankerite and
kutnohorite, respectively.
Amygdules
The amygdules are distributed highly irregularly in the
host picrite: in some parts they are missing, in foamed por-
tions of rock the total volume of amygdules is much higher
than those of rock-forming picrite matrix. The sizes are vari-
able, ranging from < 1 mm up to 1 m (Rusek & Valošek
1968). The amygdules often display zonal texture and highly
variable mineral composition precluding the construction of
an unified paragenetic sequence. Rusek & Valošek (1968)
described dolomite, ankerite, calcite, quartz, chalcedony,
opal, hematite, barite, celadonite, chlorite, goethite, rutile,
and pyrite in the amygdule fill. Recently, we have studied in
a greater detail three samples that differ in size, texture and
mineral composition.
1) Sample HH-1 consists of foamed picrite containing
~80 vol. % of white to pinkish spherical amygdules reaching
2—4 mm in diameter. All the amygdules are completely filled
up by minerals and often exhibit a zonal structure (Fig. 2a).
The dominating constituent is dolomite showing an oscillatory
growth zonality in BSE image (Fig. 2a). It contains 92—96,
3.7—7.5 and ~ 0.3 mol % of dolomite, ankerite and kutno-
horite molecules, respectively. Growth zonality of dolomite
is highlighted by tiny interpositions or inclusions of calcite
(Cal
93—98
Mag
0.7—4.3
Sid
0.1—1.8
Rdc
0.3—0.5
Str
0.1—0.3
). The youngest
central part of the amygdule is filled by either quartz (sample
HH-1) or calcite (sample HH-17).
2) Sample HH-2 is a potato-shaped amygdule 1.5 cm in
diameter, with concentric zonal texture. The oldest part is
formed by pink-grey hemispherical aggregates of dolomite
with a spherulitic arrangement of strongly elongated dolomite
individuals. Compositional zonality has not been proved using
the microprobe (Dol
94.2—94.6
Ank
5.4—5.7
Ktn
0.2—0.4
). Dolomite
aggregates are rimmed by dark brown 0.5—1 mm thick layers of
calcite I (Cal
97.8
Mag
1.2
Sid
0.4
Rdc
0.4
Str
0.2
), which are followed
by coarse-grained (up to 5 mm) rhombohedral crystals of co-
lourless to white quartz. In the central vug, white to colourless
rhombohedral crystals of calcite II are present.
3) Sample HH-3 represents the largest ( ~ 15 cm in size)
and mineralogically most interesting amygdule (Fig. 2b).
Macroscopically, three distinct types of fill can be distin-
guished. The oldest marginal part adjacent to the host rock is
composed of white to grey-pink middle-grained (0.5—1 mm)
dolomite I. The individual isometric crystals often show co-
lour zonation (milky white cores rich in fluid inclusions and
transparent rims much poorer in fluid inclusions). Microprobe
analyses showed a limited chemical variability (93—98 mol %
dolomite, 1.4—6.8 mol % ankerite, 0.2—0.6 mol % kutno-
horite). The growth zonality of the dolomite I is highlighted
by 2—3 interpositions rich in irregularly shaped patches
(Fig. 2e) of a fine-grained green phyllosilicate, essentially iso-
tropic under crossed polars. The mineral resembles celadonite
suggested by Rusek & Valošek (1968), however, with respect
to its chemical composition (Table 2) showing generally low
contents of interlayer cations (0.61—0.75 apfu K+ Na+ Ba+ Ca)
it is classified as glauconite according to Rieder et al. (1997)
– Fig. 3c. Weathered glauconite becomes brown in thin sec-
tion and loses potassium (analyse No. 6 in Table 2).
The central part of the amygdule is completely filled by a
massive fine-grained carbonate whose colour is white-to-pink
in fresh parts and brown in partly weathered parts. Moreover,
this material also fills sporadic fissures within marginal
dolomite I (Fig. 2b). In thin section, a mosaic of the very fine-
grained (5—10 µm) poorly transparent isometric carbonate
grains is observed. In BSE image, two spatially distributed
mineral phases can be distinguished. The major lighter phase
is dolomite II (Dol
86.4—88.9
Ank
11.0—13.5
Ktn
0.0—0.1
), whereas the
minor darker one is magnesite (Mag
87.5—88.6
Sid
10.1—11.2
Cal
1.0—1.2
Rdc
0.0—0.1
). The bulk chemical composition of this carbonate
mixture is given in Table 1. Based on Mg and Ca balance, the
central part of amygdule is composed of 72 and 28 (± 2) wt. %
of dolomite II and magnesite, respectively.
Spherical ball-like aggregates of greenish carbonate occur
locally in between the marginal dolomite I zone and central
fine-grained carbonate. Sometimes they are concentrated im-
mediately along the contact, sometimes they are freely
“swimming” in the fine-grained dolomite-magnesite matrix
(Fig. 2b). Each aggregate exhibits a distinct growth zonality,
visible in both optical and electron microscopes (Fig. 2c). The
crystallization started usually by compositionally homoge-
neous euhedral lenticular crystal of dolomite with lowermost
content of iron (Dol
96.8
Ank
3.1
Ktn
0.1
). An irregular grain of
barite was observed exceptionally in the centre of this dolo-
mite crystal. The most voluminous remnants of the aggre-
gate are constituted by well-elongated individuals of
carbonate oriented perpendicular to the surfaces of early do-
lomite crystals. The carbonate is mostly iron-richer
dolomite III (Ank
3.3—7.0
Ktn
0.0—0.6
), which contains tiny (up to
80 µm thick) interpositions (Fig. 2c) or layers of discrete in-
clusions of Mg-Fe-rich carbonate (Mag
38—48
Sid
38—44
Cal
13—18
Rdc
0.4—0.5
). The carbonate exhibits a detailed growth zonality
(Fig. 2d), with composition lying just around the border be-
tween Mg-siderite and Fe-magnesite in the classification
scheme of Trdlička & Hoffman (1975) – Fig. 3a. The green-
ish colour of the dolomite III is caused by spatially distributed
minute worm-like, fan-like or irregular aggregates of chlo-
rite. This mineral also occurs in the interspace matrix within
clusters of the described ball-like dolomite aggregates
(Fig. 2f). WDX analyses proved trioctahedral Fe-Mg chlo-
rites (Table 2). High Si contents (3.25—3.94 apfu) and variable
Fe/(Fe + Mg) ratios (0.49—0.66) correspond to delessite to
pennine (Melka 1965) – Fig. 3b. Finally, both the ball-like
dolomite aggregates and their interspace matrix contain rare
sulphide grains, represented by up to 60 µm large spherical
grains of older pyrite and up to 250 µm long needles of
younger millerite (Fig. 2f). Pyrite contains elevated contents
of Ni (1.3—1.9 wt. %), Co (0.2—0.4 wt. %) and Cd (0.15—
0.18 wt. %), millerite contains increased Fe (0.9—2.0 wt. %)
and Co (0.2—0.5 wt. %). The contents of Ag, Zn, As, Se, In
and Mn are below the detection limits of the microprobe.
Fluorite was rarely found to occur as irregular coarse-
grained aggregate at the contact of picrite and central fine-
331
HYDROTHERMAL MINERALIZATION IN CRETACEOUS PICRITE (WESTERN CARPATHIANS, CZECH REPUBLIC)
Fig. 2. Textural features and mineral paragenesis of hydrothermal mineralization from Hončova hůrka. a – Compositionally zonal fill of
an amygdule hosted by sample HH-1. Do – dolomite, Cc – calcite, Q – quartz. BSE image. b – Three populations of carbonate hosted
by amygdule HH-3. Early dolomite I (Do1) contains interpositions formed by glauconite (Glc) and is cut by a tiny veinlet composed of do-
lomite II+ magnesite (Do2+ Mg) filling also the central part of the amygdule. The spherical aggregates of dolomite III (Do3) are distributed
mainly along the contact of both above mentioned phases. c – Growth zonality of carbonates. The crystallization of dolomite III (Do) was
broken by interposition of Fe-rich magnesite to Mg-rich siderite (Mg—Fe). The white dots are small pyrite grains. Sample HH-3, BSE im-
age. d – Detailed growth zonality of Fe-rich magnesite to Mg-rich siderite (Mg—Fe) from figure c. e – Irregular aggregates of glauconite
(Glc) enclosed within dolomite I (Do). Sample HH-3, BSE image. f – Fine-grained aggregate of chlorite (Chl) enclosed in dolomite III
(Do) contains rounded pyrite grains (Py) which are overgrown by millerite (Mi). Sample HH-3, BSE image. All BSE images were made by
M. Dosbaba and P. Gadas.
332
DOLNÍČEK, URUBEK and KROPÁČ
grained carbonate matrix (sample PG, leg. P. Gadas). Fluo-
rite grains up to 1 cm in diameter are colourless, white and
light violet in colour.
Fluid inclusions
Fluid inclusions suitable for microthermometric analysis
were found in quartz, fluorite and carbonate from both vein
and amygdule mineralizations (Table 3).
Table 2: Representative microprobe analyses of vein glauconite (anal. No. 1—5 fresh glauconites, No. 6 weathered one) and chlorite (anal.
No. 7—10) from Hončova hůrka. The empirical formulae have been calculated for 11 and 14 anions per formula unit for glauconite and chlo-
rite, respectively.
1 2 3 4 5 6 7 8 9 10
P
2
O
5
0.06 0.08 0.02 0.05 n.d. 0.22 0.19 0.27 0.19 0.04
SiO
2
49.68 50.18 52.30 51.31 53.27 50.68 37.07 36.66 37.53 31.37
TiO
2
0.01 0.00 0.01 0.02 0.02 0.00 0.07 0.06 0.06 0.00
Al
2
O
3
8.26 8.12 9.25 9.43 7.84 12.37 9.13 9.73 9.59 14.50
Cr
2
O
3
0.00 0.01 0.00 0.00 0.00 0.03 0.02 0.00 0.02 0.04
FeO
tot
18.66 19.32 17.61 17.84 18.29 6.87 31.14 32.11 31.68 27.15
MgO
4.58 4.83 4.71 4.55 7.86 15.07 9.42 9.27 9.53 14.83
MnO
0.00 0.02 0.00 0.02 0.03 0.05 0.02 0.00 0.02 0.02
CaO
0.51 0.56 0.75 0.55 0.53 1.21 1.63 1.26 1.06 0.74
BaO
0.00 0.01 0.00 0.02 n.d. 0.03 0.00 0.05 0.00 0.01
NiO
0.07 0.06 0.02 0.08 n.d. 0.08 n.d. n.d. n.d. 0.19
ZnO
0.02 0.04 0.00 0.00 0.01 0.02 0.00 0.04 0.04 0.07
Na
2
O
0.04 0.07 0.03 0.06 0.08 0.30 0.10 0.08 0.08 0.03
K
2
O
7.46 7.09 7.19 7.14 6.35 0.96 0.17 0.14 0.13 0.01
Cl
0.04 0.04 0.04 0.05 0.03 0.09 0.02 0.01 0.00 0.02
F
0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00
Total
89.40 90.43 91.92 91.12 94.30 87.75 88.97 89.69 89.94 89.01
P
0.004 0.005 0.001 0.003
0.013 0.017 0.024 0.017 0.004
Si
3.648 3.640 3.698 3.667 3.666 3.546 3.906 3.846 3.904 3.268
Ti
0.001
0.001 0.001
0.005 0.005 0.004
Al
0.715 0.695 0.771 0.794 0.636 1.020 1.134 1.203 1.176 1.780
Cr
0.001
0.002
0.002
0.002
0.003
Fe
3+
1.146 1.172 1.041 1.066 1.053 0.402
Fe
2+
2.744
2.817
2.756
2.366
Mg
0.501 0.522 0.497 0.485 0.806 1.572 1.479 1.450 1.478 2.304
Mn
0.001
0.001 0.002 0.003 0.001
0.002 0.002
Ca
0.040 0.043 0.057 0.042 0.039 0.091 0.184 0.141 0.118 0.082
Ba
0.001
0.001
0.002
Ni
0.004 0.003 0.001 0.005
0.005
0.016
Zn
0.001
0.002
0.001
0.003
0.003
0.005
Na
0.006 0.010 0.004 0.008 0.010 0.041 0.020 0.017 0.016 0.005
K
0.699 0.656 0.649 0.651 0.558 0.085 0.022 0.019 0.017 0.001
Cl
0.005 0.005 0.005 0.006 0.004 0.011 0.003 0.002
0.004
Catsum
6.765 6.749 6.718 6.724 6.771 6.781 9.515 9.528 9.493 9.836
Sample Sample type
Mineral
FI type
Phase composition
T
h
(L+V)
T
f
T
e
T
m
ice
Salinity
HH-1
Amygdule Dolomite P
L,
L+V
62–89
–36/–45 n.d.
–1.5/–1.9
2.6–3.2
Calcite
P
L, rarely L+V
n.d.
–36/–45
–37
–1.2/–1.9
2.1–3.2
HH-2
Amygdule Quartz P
L+V,
L
56–>250
–37/–43
–34/–38
–0.3/–2.2
0.5–3.7
Calcite II
P
L, rarely L+V
n.d.
–45
n.d.
–0.2
0.4
Dolomite
P
L
–36/–42
n.d.
–0.1/–1.2
0.2–2.1
HH-3
Amygdule Dolomite
I
P
L+V
106–131
–37/–46 –35
–0.4/–1.0
0.7–1.7
Dol. II+magn
P
L+V, L
143–151
–39/–44
n.d.
–1.1/–2.1
1.9–3.5
Dolomite III
P
L+V, L
60–125
–37/–44
–37
–0.3/–1.5
0.5–2.6
PS–S?
L, L+V
60–110
–35/–41
n.d.
–0.5/–1.2
0.9–2.1
HH-7
Vein Calcite
P,
PS
L,
L+V
56–88
–35/–44
–34/–37
–0.2/–1.4
0.4–2.4
HH-17
Amygdule Calcite P
L+V,
L
110–125
–34/–43 n.d.
–0.8/–1.2
1.4–2.1
PG
Amygdule
Fluorite
P, PS
L+V, rarely L
118–172
–30/–45
n.d.
–0.2/–0.6
0.4–1.1
S
L, rarely L+V
48
–41
n.d.
–0.1
0.2
Table 3: Results of fluid inclusion microthermometry. Temperature parameters are in °C, salinity in wt. % NaCl eq.
Carbonate (calcite, dolomite) samples contain abundant
primary fluid inclusions, which are often concentrated with-
in certain growth zones, whereas other zones are essentially
inclusion-free. Such a distribution of fluid inclusions is typi-
cal especially for small dolomite amygdules in the sample
HH-1, and the early dolomite I in sample HH-3. Primary flu-
id inclusions in some other samples (fluorite, dolomite III in
sample HH-3) exhibit a regular three-dimensional distribu-
tion without apparent arrangement along growth zones.
333
HYDROTHERMAL MINERALIZATION IN CRETACEOUS PICRITE (WESTERN CARPATHIANS, CZECH REPUBLIC)
Fig. 3. Classification of some hydrothermal phases from Hončova hůrka and comparison with other teschenite/picrite hosted mineraliza-
tions. a – Carbonates of calcite group (Trdlička & Hoffman 1975). b – Chlorites (Melka 1965). c – Glauconite (Rieder et al. 1998). Fig-
ure c’ represents the lateral view on 3D diagram, figure c” is top view along the cutting plane which is marked by dashed line in figure c’.
The comparative data (outlined by dotted line) are from Urubek & Dolníček (2008), Urubek et al. (2009) and Dolníček et al. (2010).
Quartz HH-2 contains irregular clusters of fluid inclusions,
which are situated mostly around the centre of the crystal,
whereas the rims are essentially inclusion-free. Very rare
solitary fluid inclusions have been observed in fine-grained
dolomite II + magnesite mixture in sample HH-3 as well as in
late calcite crystal in sample HH-2.
Most primary inclusions show equant shapes and sizes rang-
ing between 5 and 26 µm. All samples contain aqueous
inclusions only. At room temperature, a coexistence of one-
phase (L-only) and two-phase (L + V) inclusions with essen-
tially constant liquid-vapour ratios can be observed in most
samples (the volume of gaseous phase does not exceed ca. 5 %).
Both types of fluid inclusions occur together within the same
structures (clusters, growth zones, etc.), however, the abun-
dance of both compositional types varies significantly between
samples (Table 3). The absence of vapour bubbles is often ob-
served in smaller ( < 10 µm) inclusions implying that metasta-
bility of the bubble nucleation could play a role. Few samples
contain exclusively the L-only inclusions including those of
larger sizes (e.g. calcite HH-2, dolomite HH-2) suggesting very
low trapping temperatures (below ca. 50 °C; Goldstein &
Reynolds 1994). Irregularly shaped inclusions in quartz HH-2
exhibit a wider variability in liquid-vapour ratios (up to ~ 30 %
of vapour phase), possibly caused by necking-down.
Homogenization temperatures of primary L+ V inclusions
in carbonates and fluorite range between 56 and 172 °C
(Fig. 4a). Quartz HH-2 alone exhibit even wider spread be-
tween 56 and > 250 °C consistent with necked nature of fluid
inclusions. In all samples, the inclusions freeze at tempera-
tures from —34 to —45 °C. The vapour bubble was often elim-
inated by expanding ice, and during subsequent heating a
metastable ice melting is common. In such cases, the fluid
inclusions have been artificially stretched by overheating to
~ 200—250 °C prior to cryometric runs. The rarely observed
eutectic temperatures are from —35 to —38 °C, suggesting the
NaCl-MgCl
2
-H
2
O fluid composition. The last ice melts at
temperatures between —0.2 and —2.2 °C (Fig. 4b) indicating
bulk fluid salinities between 0.4 and 3.7 wt. % NaCl eq.
(Bodnar 1993). There are no systematic differences between
L-only and L+ V fluid inclusions in their cryometric
334
DOLNÍČEK, URUBEK and KROPÁČ
parameters, which support the idea about a metastable nature
of the liquid inclusions.
The irregularly-shaped secondary fluid inclusions ar-
ranged along healed microfractures are sometimes present in
studied samples. They have generally similar microthermo-
metric parameters as the primary fluid inclusions (Table 3).
C, O, S and Sr isotopes
Eight samples of carbonates were analysed for carbon and
oxygen isotope compositions (Table 4). The
δ
18
O values
vary between —0.6 and —6.5 ‰ PDB, and the
δ
13
C values are
between —4.5 and + 1.6 ‰ PDB.
One sample of sulphide sulphur, extracted from bulk
amygdule HH-3, showed the
δ
34
S value as high as + 1.1 ‰
CDT. Two attempts to collect a sufficient quantity of sulphur
compounds for isotope analysis from host rocks were not
successful.
Fig. 4. Results of microthermometry of primary and primary-sec-
ondary fluid inclusions from Hončova hůrka. a – Histogram of ho-
mogenization temperatures of L + V inclusions. b – Histogram of
temperatures of last ice melting. c – T
h
—T
m
ice plot. n – number of
measurements.
The
87
Sr/
86
Sr ratios were determined in five carbonate sam-
ples and two host rocks (Table 5). The lowest present-day
87
Sr/
86
Sr ratios show both picrite samples (0.7045 and 0.7048
for the least and the most altered sample, respectively). All the
hydrothermal carbonates exhibit considerably higher
87
Sr/
86
Sr
ratios between 0.7060 and 0.7068. The low values character-
ize the samples from picrite host rocks, the highest one
originated from a vein calcite hosted by claystone.
Trace elements
Five carbonate separates and two samples of host picrite
have been analysed for trace elements (Table 6). Low contents
of elements incompatible with calcite structure (e.g. Ga, Zr,
Rb) indicate a negligible contamination of carbonates by the
host rock and/or hydrothermal silicate mineral phases. Low
contents of heavy metals (except for Ni), low to moderate con-
tents of Ba (32—515 ppm) and moderate to high contents of Sr
(400—3700 ppm) were found in calcite and dolomite samples.
The Ni content is elevated in all picrite-hosted samples
(between 23 and 86 ppm), whereas the claystone-hosted calcite
contains only 1.5 ppm Ni.
The two host rock samples exhibit similar chondrite-nor-
malized REE patterns although they differ in both degree of
alteration and total REE concentrations (164 and 281 ppm
for least altered and strongly altered picrite, respectively).
Both samples show LREE-enriched patterns (La
N
/Yb
N
= 20
and 34) without any Ce and Eu anomalies (Fig. 5). In con-
trast, data on hydrothermal carbonates can be divided into
two subsets differing in both absolute amount of REE and
shape of normalized REE patterns. Group 1 is characterized
by high content of REE (68—216 ppm) and LREE enrich-
ment, which are both comparable to those of host picrite.
The Ce and Eu anomalies are missing. Group 2 contains
samples showing very low REE contents (0.12—0.93 ppm),
with many elements below detection limit. The chondrite-
normalized patterns are much more well-balanced in terms
of LREE/HREE ratio (Fig. 5). No positive Eu anomalies are
present but two of three samples exhibit a negative Ce anom-
aly (Ce/Ce* = 0.68 and 0.72).
Fig. 5. REE chondrite-normalized patterns of hydrothermal carbon-
ates and host picrite. Normalization values are from Anders &
Grevesse (1989).
335
HYDROTHERMAL MINERALIZATION IN CRETACEOUS PICRITE (WESTERN CARPATHIANS, CZECH REPUBLIC)
Table 4: Carbon and oxygen isotope composition of carbonates and
δ
13
C and
δ
18
O values of their parent fluids calculated for the given
temperature.
Table 5: Rb—Sr elemental and isotope data on samples from Hončova hůrka and calculated initial
87
Sr/
86
Sr
i
ratios for age of 120 Ma. The
Rb contents in parentheses represent the maximum possible value, for which the initial
87
Sr/
86
Sr
i
ratio was calculated.
Table 6: Trace element abundances in carbonate and picrite samples from Hončova hůrka. All values are in ppm. The contents of Ag, As,
Au, Be, Bi, Cd, Cs, Hg, Sb, Se, Sn, Tl and W in all samples were below or just around the detection limits ranging between 0.1 and 1 ppm.
Sample Sample
type Mineral
13
C carb.
(‰ PDB)
18
O carb.
(‰ PDB)
18
O carb.
(‰ SMOW)
Temperature
(°C)
13
C fluid
(‰ PDB)
18
O fluid
(‰ SMOW)
HH-1
Amygdule
Dolomite –3.7 –1.4 29.4 62–89
–6.1/–6.8
+2.9/+7.0
HH-2
Amygdule
Dolomite –2.2 –0.6 30.3 30–50
–5.8/–7.0
–2.6/+1.6
Calcite
II
–4.5
–5.8
24.9
30–50
–7.9/–9.1
–2.0/+1.7
HH-3
Amygdule Dolomite
I
–3.4
–0.6
30.3
106–131 –5.6/–6.5
+10.0/+12.7
Dol.
II+magn
+1.6
–0.7
30.2
143–151
–0.5/–0.9
+13.7/+14.4
Dolomite
III
–3.8
–0.8
30.1
60–125
–5.9/–6.9
+3.2/+11.9
HH-5
Vein Dolomite –0.6 –5.0 25.7 n.d. n.d. n.d.
HH-7
Vein Calcite –3.5 –6.5 24.2 56–88
–5.7/–6.6
+1.9/+6.3
Sample
Sample type
Mineral
Rb (ppm)
Sr (ppm)
87
Sr/
86
Sr
m
±2
87
Sr/
86
Sr
i
HH-1
Amygdule Dolomite
1.0 1590 0.706586 0.000017 0.706583
HH-3
Amygdule
Dolomite I
(0.4)
400
0.706521
0.000020
0.706516
Dol.
II+magn
(0.4)
3700
0.706000
0.000018
0.705999
HH-5
Vein Dolomite (0.4)
800
0.705968
0.000019
0.705966
HH-7
Vein Calcite
(0.4)
3390
0.706806
0.000019
0.706805
HH-1
Host rock
Picrite
34.7
900
0.704780
0.000019
0.704590
HH-4
Host rock
Picrite
34.4
510
0.704503
0.000018
0.704170
Sample
HH-4 HH-1 HH-1 HH-3 HH-3 HH-5 HH-7
Mineral
Picrite
Picrite Dolomite Dolomite
I Dol.
II+magn Dolomite Calcite
Sample type
Host rock
Host rock
Amygdule
Amygdule
Amygdule
Vein
Vein
Ba
606
1150
515
262
86.9
74.7
32.1
Co
95.6
50.7
66.7
1.4
4.1
0.5
<0.5
Cu
37.8
50.1
16.7
0.5
0.1
0.2
0.4
Ga
18.9
13.6
<0.5
<0.5
<0.5
<0.5
<0.5
Hf
3.7
4.4
<0.5
<0.5
<0.5
<0.5
<0.5
Mo
1.9
0.2
<0.1
0.5
0.2
0.8
<0.1
Nb
61.9
86.6
<0.5
<0.5
<0.5
<0.5
<0.5
Ni
616
307
49.0
22.9
86.2
24.0
1.5
Pb
1.8
1.5
1.4
0.1
<0.1
0.6
0.7
Rb
34.4
34.7
1.0
<0.5
<0.5
<0.5
<0.5
Sc
19
28
n.d.
n.d.
n.d.
n.d.
n.d.
Sr
510
900 1590
400 3700
800 3390
Ta
3.2
4.7
<0.1
<0.1
<0.1
<0.1
<0.1
Th
4.2
9.2
0.9
<0.1
0.1
<0.1
0.2
U
1.8
2.3
0.8
1.9
1.4
0.3
<0.1
V
200
250
39
7
16
6
<5
Zn
72
51
45
3
7
4
<1
Zr
154
174
1.9
7.9
6.3
1.5
<0.5
Y
18.3
22.6
20.2
0.1
0.3
0.4
10.4
La
36.3
70.3
52.9
0.4
0.1
0.3
12.4
Ce
66.7
117
91.9
0.4
<0.1
0.4
26.6
Pr
8.15
13.71
10.34
0.05
<0.02
0.06
3.44
Nd
32.0
51.7
38.2
<0.3
<0.3
<0.3
15.0
Sm
5.59
8.78
6.23
<0.05
<0.05
<0.05
3.12
Eu
1.81
2.76
2.05
<0.02
<0.02
<0.02
0.98
Gd
4.86
6.76
5.64
0.08
<0.05
<0.05
3.03
Tb
0.72
1.00
0.86
<0.01
0.01
0.01
0.46
Dy
3.45
4.90
3.86
<0.05
<0.05
0.07
1.88
Ho
0.66
0.78
0.61
<0.02
<0.02
<0.02
0.29
Er
1.60
1.85
1.50
<0.03
<0.03
<0.03
0.61
Tm
0.24
0.28
0.19
<0.01
0.01
0.01
0.08
Yb
1.27
1.38
1.17
<0.05
<0.05
<0.05
0.42
Lu
0.19
0.19
0.16
<0.01
<0.01
0.01
0.06
REE
164
281
216
0.93
0.12
0.86
68.4
La
N
/Yb
N
19.8
35.3
31.4
n.d.
n.d.
n.d.
20.5
Ce/Ce*
0.93
0.90
0.94
0.68
n.d.
0.72
0.98
Eu/Eu*
1.06
1.09
1.05
n.d.
n.d.
n.d.
0.97
336
DOLNÍČEK, URUBEK and KROPÁČ
Discussion
Source of hydrothermal fluids
The available geochemical data offer several lines of evi-
dence that the major source of parent hydrothermal fluids
was not related to the host picrite:
Salinity of hydrothermal fluid. It is well documented that
silicic magma can exsolve hydrosaline fluid phase during crys-
tallization (e.g. Burnham 1979; Cline & Bodnar 1991; Webster
2004). The actual fluid salinity depends on (i) pressure (i.e.
depth of intrusion) and (ii) progress in magmatic crystallization
(Cline & Bodnar 1991; Fall et al. 2007). At pressures about
0.5 kbars (i.e. in shallow settings), the salinity of the earliest
magmatic fluid is low and it increases rapidly during crystalliza-
tion. At pressures about 1.3 kbars the salinity shows only minor
changes towards higher values at the end of crystallization. In
deep (pressures higher than ca. 2 kbars) systems the first fluid
exsolved from the melt has high salinity and salinity decreases
as crystallization proceeds. If the fluids at Hončova hůrka were
magmatic in origin, then the observed occurrence of low-sa-
linity solutions (that have to operate during the final stage of
magmatic crystallization) would indicate the deep-seated high-
pressure conditions (in depth at least 20 km under hydrostatic
conditions). However, the occurrence of carbonate-rich sedi-
ments in the surroundings of the picrite body indicates that the
seafloor was situated well above CCD at Hončova hůrka site
( ~ 4—5 km at maximum, cf. van Andel Tjeerd 1975 in Kiss et al.
2008). Thus the magmatic model alone is incompatible with
observed data and geological situation, and alternative fluid
source(s) must have been involved. Mixing of two fluid end-
members which differ in temperature and salinity can be in-
ferred from the T
h
-T
m
ice diagram for all carbonate samples
(Fig. 4c). The higher-salinity fluid endmember could possibly
be seawater as is indicated from salinity reaching up to
~3.5 wt. % and the Na-Mg-Cl salt composition. The origin of
the low-salinity fluid endmember cannot be specified from
salinity data alone, but meteoric water as the usually most avail-
able source could be safely rejected in our submarine setting.
δ
18
O and
δ
13
C of the hydrothermal fluid. The fluid
δ
18
O
and
δ
13
C characteristics have been calculated from mineral
δ
18
O and
δ
13
C data using the measured homogenization tem-
peratures or estimated crystallization temperatures (Table 4).
It should be noted that the use of pressure-uncorrected T
h
val-
ues leads to underestimated fluid
δ
18
O and
δ
13
C values. Fortu-
nately the low-pressure systems need only insignificant
correction. At Hončova hůrka the pressure does not exceed
500 bars ( = depth of 5 km under hydrostatic conditions). A
more detailed quantification is not possible from the available
data. If the pressure attained the maximum value of 500 bars,
then the measured T
h
values need to be shifted by 15—25 °C to
obtain real trapping temperatures (Flincor software; Brown
1989). Such an increase of temperature shifts both the fluid
δ
18
O and
δ
13
C values by 2.0—2.5 and 0.5—1.5 ‰, respectively.
These (maximum possible) shifts do not affect the interpreta-
tion of carbon and water sources indicated below.
The fluid
δ
18
O values are widely scattered between ~ 0 and
~+14 ‰ SMOW overlapping significantly the range of
magmatic waters ( + 5 to + 10 ‰ SMOW; Sheppard 1986).
The observed variability of
δ
18
O values could be explained
by 1) highly variable exchange of oxygen between rocks and
fluid phase, or 2) mixing of two (or more) fluids with con-
trasting isotope compositions (e.g. seawater with a near-zero
δ
18
O value could mix with metamorphic, diagenetic or or-
ganic waters with highly positive
δ
18
O values; cf. Sheppard
1986). In the given geological setting the generation of “di-
agenetic” low-salinity waters has been documented in clay-
rich sedimentary sequences (cf. Polách et al. 2008; Dolníček
& Polách 2009). The dewatering of clays could have been
associated with burial compaction and/or heating-up caused
by intrusion of teschenite/picrite magma.
Most samples show a limited range of calculated fluid
δ
13
C values (—6 to —9 ‰ PDB; Table 4) which largely over-
lap the values typical of “magmatic” carbon (—5 to —8 ‰
PDB; Hoefs 1997). However, the same range exhibits also
“carbon of the homogenized Earth’s crust”, averaged from
various crustal sources during fluid evolution. Nevertheless
the influence of external sources is manifested by near-zero
fluid
δ
13
C value of dolomite II from sample HH-3 indicating
the source of carbon in sedimentary carbonates (limestones).
REE signature of hydrothermal fluid. The REE data on hy-
drothermal carbonates cluster in two internally consistent
groups. Such a behaviour does not favour the interpretation
in terms of crystallization from a single fluid enriched in
REE-complexing ligands, where a continuous change in
REE patterns of precipitating carbonates can be expected
during crystallization. Rather, our data could reflect different
sources of REE in the hydrothermal fluids and/or interaction
with different rocks. While Group 1 samples could have
been inferred from host picrite, Group 2 shows REE concen-
trations and patterns typical of Cretaceous marine limestones
(Bellanca et al. 1997; Temur et al. 2009). The negative Ce
anomaly in Group 2 samples could reflect the presence of
seawater in the hydrothermal system (McLennan 1989); or a
signature inherited from limestones.
Fig. 6.
87
Sr/
86
Sr
i
vs. 1000/Sr plot for carbonate and picrite samples
from Hončova hůrka.
337
HYDROTHERMAL MINERALIZATION IN CRETACEOUS PICRITE (WESTERN CARPATHIANS, CZECH REPUBLIC)
Sr isotope composition of hydrothermal fluid. Table 5 gives
the “initial”
87
Sr/
86
Sr
i
ratios calculated for age of 120 Ma.
This age represents the solidification age of host igneous rock,
and the uppermost possible age of the studied mineralization.
Due to extremely low Rb/Sr ratios, the
87
Sr/
86
Sr
i
ratios of the
carbonate samples are (within the analytical uncertainty)
indistinguishable from the present-day
87
Sr/
86
Sr values. The
87
Sr/
86
Sr
i
ratios of all hydrothermal carbonates are much
higher (0.7060 to 0.7068) then those of host picrite (0.7042
and 0.7046) thus indicating an external source of at least a
part of the strontium in the hydrothermal fluid. The possible
sources of more radiogenic strontium are (i) Lower
Cretaceous seawater (
87
Sr/
86
Sr = 0.7071 to 0.7075; Veizer et
al. 1999), (ii) marine limestones occurring in rock sequences
of the Silesian Unit, and/or (iii) siliciclastic sedimentary rocks.
The probable influence of more then two sources of stron-
tium is indicated by a lack of correlation between
87
Sr/
86
Sr
and 1000/Sr (Fig. 6).
Evidence for interaction with host picrite
As indicated above, many mineral-forming elements and
major part of fluid phase have been derived out of the picrite
body. However, the picrite body served not only as a suitable
conduit for fluid flow and site for the associated precipitation
of mineral phases, but there is evidence that the igneous rock
has also actively interacted with hydrothermal fluids.
Picrite alteration and the nature of hydrothermal mineral-
ization. The nature and mineral composition of secondary
phases in host picrite are similar to those in the studied
amygdule and vein parageneses: in both cases phyllosilicates
and carbonates predominate. Therefore, it is probable that
pervasive alteration of picrite was coeval with hydrothermal
activity giving rise to the open-space mineralization (cf.
Kudělásková et al. 1990). Two specific features can further
support this opinion. 1) The vein and amygdule mineraliza-
tion is characterized by a pronounced enrichment in magne-
sium. The iron-poor dolomite as a dominating gangue
mineral as well as the presence of magnesite is not known
from other sites in the sub-Beskydy magmatic province. The
occurrence of these phases is coupled with an increase of
FeO/(FeO + MgO) ratio from 0.35 to 0.43 during increasing
intensity of picrite alteration (Table 1) implying that magne-
sium was preferentially leached from the rock by circulating
fluids. 2) The mineralization at Hončova hůrka contains ele-
vated contents of Ni in pyrite and carbonates. In addition, a
trace amount of millerite is described for the first time here.
The Ni-enrichment of the fluid phase is coupled with Ni-de-
pletion of altered rocks (Table 6), again suggesting a source
of this element in the host rock (olivine in the sample HH-4
contains up to 0.23 wt. % NiO; unpubl. microprobe data of
the authors). Both the above mentioned findings are unique
for the Hončova hůrka site only. This can evoke an idea
about low water-rock ratio during alteration/mineralization
process in order to allow the increasing Mg and Ni concen-
tration in the fluid to the level at which the Mg- and Ni-rich
phases can precipitate. However the available data on fluid
salinity and REE fractionation from Hončova hůrka are com-
parable to those from other (teschenite/picrite-hosted and do-
lomite-magnesite-millerite free) mineralizations (cf. Polách
2008; Urubek 2009) suggesting that hydration reactions did
not shift the contents of both salts and strong REE-complex-
ing ligands in the remaining fluid phase. This behaviour
would favour the open-system fluid circulation and higher
water-rock ratios (cf. Dolníček et al. 2010; Kiss et al. 2008).
δ
34
S of the hydrothermal fluid. The near-zero
δ
34
S value of
fluid-precipitated sulphides is compatible with mantle-derived
sulphur (Hoefs 1997), which could have been leached from
the host picrite. However, it must be noted that such a
δ
34
S
value can also be generated by a variety of other processes.
Conclusion
The vein and amygdule mineral assemblages hosted by a
submarine effusive body of Lower Cretaceous picrite at
Hončova hůrka Hill exhibit largely similar fluid inclusion
and geochemical parameters although they can differ signifi-
cantly in mineral compositions and texture. This suggests that
they are related to the same mineralizing event. Besides the
dominating dolomite, magnesite, siderite, quartz, calcite, chlo-
rite, glauconite, fluorite, barite, pyrite and millerite were also
identified. The parent fluids are characterized by variable but
generally low temperatures (< 50—170 °C), low salinities (0.4
to 3.7 wt. %), low content of strong REE-complexing ligands,
near-zero to highly positive (up to + 14 ‰ SMOW)
δ
18
O val-
ues, near-zero to negative (down to —9 ‰ PDB)
δ
13
C values,
near-zero
δ
34
S value, and initial
87
Sr/
86
Sr ratios much more ra-
diogenic (0.7060 to 0.7068) than those of the host picrite
(0.7042 and 0.7046). These data are inconsistent with an
orthomagmatic picrite-related origin of the parent fluids,
which are interpreted to be predominantly of external origin
derived by mixing of seawater with diagenetic waters, the lat-
ter produced by dewatering of clay minerals in associated
flysch sediments. The isotope and REE signatures indicate
interaction of at least a part of the fluids with sedimentary
carbonates prior to mineral precipitation. The interaction of
the fluids with the host picrite led to strong alteration of the
rock and leaching of some elements (Mg, Ni, S, partly also
REE) that participated later during precipitation of vein- and
amygdule-hosted hydrothermal phases.
Acknowledgments: The study was supported by Project
GAČR 205/07/P130. M. Dosbaba and P. Gadas (MU Brno)
are thanked for assistance during microprobe work. Special
thank is to P. Gadas for providing the fluorite sample. The
isotope analyses conducted by I. Jačková, Z. Lněničková,
V. Janoušek and V. Erban (ČGS Praha) are highly appreciat-
ed. Constructive comments by three journal reviewers
(V. Hurai, F. Molnár and an anonymous one) helped to
improve the initial draft of the manuscript.
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